Accepted Manuscript
Late Ediacaran-Cambrian structures and their reactivation during the Variscan and Alpine cycles in the Anti-Atlas (Morocco)
A. Soulaimani, A. Michard, H. Ouanaimi, L. Baidder, Y. Raddi, O. Saddiqi, E.C. Rjimati
PII: S1464-343X(14)00131-9
DOI: http://dx.doi.org/10.1016/j.jafrearsci.2014.04.025
Reference: AES 2037
To appear in: African Earth Sciences Received Date: 15 November 2013 Revised Date: 8 April 2014 Accepted Date: 25 April 2014
Please cite this article as: Soulaimani, A., Michard, A., Ouanaimi, H., Baidder, L., Raddi, Y., Saddiqi, O., Rjimati, E.C., Late Ediacaran-Cambrian structures and their reactivation during the Variscan and Alpine cycles in the Anti- Atlas (Morocco), African Earth Sciences (2014), doi: http://dx.doi.org/10.1016/j.jafrearsci.2014.04.025
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Late Ediacaran-Cambrian structures and their reactivation during the 1
Variscan and Alpine cycles in the Anti-Atlas (Morocco) 2
3
A. Soulaimania; A. Michardb, H. Ouanaimic; L. Baidderd; Y. Raddie; O. Saddiqid, E.C. Rjimatie
4
5
a Department of Geology, Faculty of Sciences-Semlalia, Cadi Ayyad University, P.O. Box 2390, Marrakech, Morocco
6
b 10, rue des Jeûneurs, 75002 Paris, France
7
c Departement of Geology, ENS, Cadi Ayyad University, BP S2400, Marrakech, Morocco
8
d Hassan II University, Faculty of Sciences Aïn Chock, Lab. of Géosciences, BP 5366 Maârif, Casablanca, Morocco
9
f Direction du Développement Minier, Ministère de l'Energie et des Mines, B.P. 6208, Rabat Instituts, Haut Agdal, Rabat, Maroc
10 11 12
Abstract: The post-Pan-African evolution of the northern border of the West African Craton is
13
largely controlled by the remobilisation of Late Neoproterozoic basement faults. The Upper
14
Ediacaran volcanic and volcano-sedimentary sequences of the Ouarzazate Group show dramatic and
15
rapid thickness changes, consistent with active extensional faulting associated with post-orogenic
16
collapse and incipient continental rifting. The geometry and kinematics of these faults differ from
17
west to east in the Anti-Atlas. N- to NE-trending faults dominate in western Anti-Atlas in response
18
to E-W to NW-SE pure extension, while a transtensive opening regime characterize the central (Bou
19
Azzer) and eastern (Saghro-Ougnate) Anti-Atlas.
20
The marine incursion in the west-central Anti-Atlas during the late Ediacaran-Early Cambrian
21
occurred without major geodynamical break between the continental Ouarzazate Group and marine
22
sediments of the Adoudou Fm. Extensional tectonics went on during the Early Cambrian, being
23
concentrated in the western and central parts of the belt. From Middle Cambrian to Lower Devonian
24
and mainly due to thermal subsidence, the Anti-Atlas basement was buried under marine sediments
25
with dominant south-derived detrital input. Basement faults control the distribution of subsiding
26
versus shallow areas. During the Middle-Late Devonian, the dislocation of the Saharan platform
27
occurred, mainly in the eastern Anti-Atlas where Precambrian faults were also remobilized during
28
the Early Carboniferous.
29
During the Variscan orogeny, the Paleozoic series of the Anti-Atlas basin were involved in
30
folding tectonics, concomitant with the uplift of Proterozoic basement blocks bounded by inherited
31
basement faults. The pre-existing rift-related faults were variably inverted across the Anti-Atlas. In
32
the westernmost part of the belt, Variscan shortening induced positive inversions along the
33
remobilized basement faults, but in some cases, some faults preserved an apparently normal throw.
34
Some hidden basement faults accommodate the Variscan shortening by strike-slip movement
35
expressed by en echelon fold pattern in the overlying cover.
36
The Mesozoic-Cenozoic evolution of the Anti-Atlas is again marked by the reactivation of the
37
basement faults. During the Triassic, the Anti-Atlas belongs to the uplifted shoulder of the Atlasic-
38
Atlantic rift zone. Remobilisations of paleofaults again play a significant role in the weak burial of
39
the Anti-Atlas during the Late Cretaceous-Eocene before the Neogene exhumation related to the
40
Africa-Europe collision. Hence the structural evolution of the Anti-Atlas during the Paleozoic to
41
Present times has been heavily dependent on the fault pattern inherited from the Late Ediacaran-
42
Cambrian rifting evolution at the northern fringe of the WAC, already deeply affected by the Pan-
43
African orogeny.
44 45
Keywords: Anti-Atlas, Basement faults, Reactivation, Rifting, Inversion, Neoproterozoic,
46
Paleozoic, Variscan orogeny, Atlas orogeny
47
1. Introduction
48 49
The processes of fault reactivation and tectonic inversion of pre-existing faults are among
50
the most important mechanisms in continental tectonics (Coward, 1994; Holdsworth et al., 2001).
51
During superimposed orogenic cycles, continental deformation is concentrated in large-scale fault
52
networks that are frequently reactivated over long time scales. The Anti-Atlas Mountains of sub-
53
Saharan Morocco, here considered, extends at the northern border of the West African Craton
54
(WAC; Fig. 1, insert). The WAC formed 2 Ga ago through the Eburnean continental collage
55
against the Archean nucleus of north-western Africa (Rocci et al. 1991; Schofield et al., 2006).
56
The Anti-Atlas basement includes remnants of Paleoproterozoic continental crust similar to the
57
Eburnean crust of the northern part of the Reguibat Shield (Choubert, 1963; Aït Malek et al.,
58
1998; Thomas et al., 2002; Walsh et al., 2002; Gasquet et al., 2004, 2005, 2008). However, the
59
Paleoproterozoic basement of the Anti-Atlas display evidence of multiscale remobilisations
60
related to three subsequent orogenic cycles, namely the Pan-African, Variscan and Alpine cycles,
61
dated from the Neoproterozoic, Paleozoic and Mesozoic-Cenozoic, respectively. So, the Anti-
62
Atlas mountain range offers numerous opportunities to analyse the mechanisms of structural
63
heritage and tectonic inversion.
64
After an apparent long-lived Mesoproterozoic quiescence only recorded by some sediments
65
south of the Reguibat Shield (Rooney et al., 2010) and by mafic dykes emplacement in the
66
western Anti-Atlas (El Bahat et al., 2013), the Pan-African cycle (900-550 Ma) began with an
67
important dislocation of the northern and eastern margins of the WAC, related to the break-up of
68
the Rodinia supercontinent (Li et al., 2008). In the south-western part of the Anti-Atlas (from the
69
Kerdous to Zenaga inliers; Fig. 1), the dislocation of the Taghdout-Lkest Group platform and its
70
intrusion by mafic rocks were concomitant with the formation of oceanic domains and volcanic
71
arcs further in the north (Leblanc, 1975; Saquaque et al., 1989; Hefferan et al., 2000; Admou and
72
Juteau, 2000; Gasquet et al., 2005, 2008). The subsequent Pan-African compressional events with
73
a climax around 650 Ma are responsible for the accretion of island-arc and oceanic crust remnants
74
to the northern edge of the WAC. These oceanic allochthons are exposed in the Siroua and Bou
75
Azzer inliers (Leblanc and Moussine-Pouchkine, 1994; Saquaque et al., 1989; Hefferan et al.,
76
2000; Thomas et al., 2002, 2004; Inglis et al., 2004; D’Lemos et al., 2006; Soulaimani et al.,
77
2006; Walsh et al., 2012). Pan-African thrusts are prevalent in the central Anti-Atlas along the
78
Anti-Atlas Major Fault (AAMF; Choubert, 1947) whereas elsewhere in the Anti-Atlas, Pan-
79
African overprint mainly corresponds to local reactivations of Eburnean fractures along major
80
shear zones (Hassenforder, 1987). Although the AAMF is an important fracture zone, having
81
allowed the Cryogenian oceanic complex to be preserved, it does not mark the northern boundary
82
of the metacratonic border of the WAC, which extends beneath the Saghro and Ougnat inliers of
83
eastern Anti-Atlas (Ennih and Liégeois, 2008).
84
In the waning stages of the Pan-African collision, i.e. during the Ediacaran (Fig. 2), the
85
northern margin of the WAC experienced seemingly transpressional mild deformation (Gasquet
86
et al., 2008), then post-orogenic collapse, tilting of basement blocks and the creation of
87
continental basins filled by the, upper Ediacaran volcaniclastic series of the Ouarzazate Group.
88
This post-orogenic context changed into rifting during the latest Ediacaran- Lower Cambrian
89
(Taroudant and Tata Groups, from bottom to top) as testified by the interleaved alkaline basalts
90
and conspicuous synsedimentary deformations (Soulaimani et al., 2003; Buggisch and Siegert,
91
1988; Algouti et al., 2002; Benssaou and Hamoumi, 2003; Saddiqi et al., 2011). Later in the
92
Phanerozoic, the inherited basement fractures continue to play a major role in the tectonic
93
evolution of the Anti-Atlas. The Ediacaran-Early Cambrian extensional faults have a strong
94
influence on the Cambrian-to Early Carboniferous sedimentation and ensuing Variscan structures
95
that developed in the course of the Late Carboniferous (Piqué et al., 1987; Soulaimani, 1998;
96
Michard et al., 2010). At that time, the Anti-Atlas area is the foreland fold belt of the Variscan
97
Orogen that extended in the Meseta Block to the north and the Mauritanides to the southwest. The
98
role of the Precambrian basement faults in the Paleozoic sedimentation and folding has been
99
documented in several studies (Jeannette and Piqué, 1981; Soulaimani et al., 1997; Raddi et al.,
100
2007; Soulaimani and Burkhard, 2008; Michard et al., 2010). Conspicuous variation of the
101
Variscan fold trends along the Anti-Atlas have been interpreted as the result of basement faults
102
reactivation, particularly around the basement uplifts in the inherited zones of crustal weakness
103
(Leblanc, 1972, 1975; Donzeau, 1974; Jeannette and Piqué, 1981; Hassenforder, 1987;
104
Soulaimani, 1998; Belfoul et al., 2001).
105
By the end of the Variscan orogeny, most of the Anti-Atlas structure was established.
106
Nevertheless, part of the inherited basement faults were reactivated again during .the Mesozoic-
107
Cenozoic evolution of the Anti-Atlas. This must be taken into account to understand the present
108
day topography of this rejuvenated belt, next and parallel to the coeval High Atlas. Recent studies
109
based on thermochronology on apatite clarified the rate and timing of the Anti-Atlas vertical
110
movements during post-Variscan times (Malusà et al., 2007; Ruiz et al., 2010, Oukassou et al.,
111
2013).
112
The present paper explores the location and importance of the upper Ediacaran basement
113
faults and their further reactivation in response to continental extension and compression. We
114
propose a description of some typical Anti-Atlas structures during the successive tectonic events
115
in the light of new field observations and of apatite thermochronology. Our main purpose is to
116
point out the permanent tectonic heritage from the late Neoproterozoic (upper Ediacaran) to
117
Present.
118 119
2. Geological setting
120
121
The Anti-Atlas mountain range (Fig. 1) displays a N60E-striking axis where Precambrian
122
rocks crop out as extended antiformal inliers (“boutonnières”) surrounded by Cambrian marine
123
series. The folded Ordovician-Early Carboniferous series of the Anti-Atlas extend widely south of
124
the mountain axis up to the undeformed part of the WAC cover, i.e. the Tindouf Basin and
125
overlying Cenozoic plateaus (“hamadas”). In the north, the Anti-Atlas is separated from the High
126
Atlas by the South Atlas Fault or South Atlas Front (SAF) and by narrow, discontinuous Neogene
127
foreland basins, namely the Souss and Ouarzazate basins. Between the two basins, the Pan-
128
African basement of the Marrakech High Atlas lies in direct contact with the basement of the
129
Anti-Atlas cropping out in the Siroua Plateau (dominated by the Neogene Siroua volcano)
130
through the steep northward-dipping SAF.
131
The Anti-Atlas mountain range is divided into two contrasting parts by the Anti-Atlas Major
132
Fault (AAMF; Choubert, 1947). South of the AAMF, Paleoproterozoic schists and granites form
133
a large part of the inliers with granite intrusions dated at ca. 2 Ga (Aït Malek et al., 1998; Walsh
134
et al., 2002; Thomas et al., 2002; Gasquet et al., 2005; O'Connor et al.,2010; Hafid et al. 2013;
135
Soulaimani et al., 2013). These Eburnean rocks are overlain by remnants of their lower
136
Neoproterozoic, shallow-water cover series (Jbel Lkest-Taghdout Group) folded and
137
recrystallized during the Pan-African orogeny. Contrastingly, within the AAMF corridor and
138
immediately north of it (Siroua and Bou Azzer inliers), the metamorphic rock units overlain by
139
the Ediacaran volcaniclastic groups are either Neoproterozoic ophiolites or coeval arc-related
140
gneiss and plutons accreted to the northern edge of the WAC. Their obduction as tectonic slices
141
along the Bou Azzer-Siroua suture occurred during two main Pan-African events at 760 Ma
142
(D’Lemos et al., 2006) and 650 Ma (Saquaque et al., 1989; Hefferan et al., 2000; Thomas et al.,
143
2002; El Hadi et al., 2010).
144
The late Pan-African, post-paroxysmal Saghro and Bou Salda Groups (lower Ediacaran) crops
145
out mainly in the eastern and central parts of the Anti-Atlas, whereas the upper Ediacaran
146
Ouarzazate Group extends all over the belt and occupies more than half of the total inlier surface
147
(Fig. 1). The conglomerates and volcanics (mainly andesites and rhyolites) of the latter group are
148
associated with high-K, calc-alkaline plutons dated at 580-550 Ma (Thomas et al., 2002, 2004;
149
Inglis et al., 2004; Levresse et al., 2004; Gasquet et al., 2005).
150
The Ouarzazate Group formations are, as a rule, conformably overlain by the uppermost
151
Ediacaran-Lower Cambrian carbonate deposits (Taroudant and Tata Groups, from bottom to top),
152
although an unconformity is observed by place (see below, sect 3.2). The almost continuous
153
Cambrian-Lower Carboniferous series, ~6 to 10 km-thick, are deformed into conspicuous fold
154
trains (Fig. 1), upright and generally open in most areas or reclined and associated with thrusts in
155
the westernmost regions (Soulaimani, 1998; Belfoul et al., 2001; Helg et al., 2004; Soulaimani
156
and Burkhard, 2008; Raddi et al., 2007; Michard et al., 2008, 2010). This Late Carboniferous,
157
strong Variscan folding was accompanied by weak recrystallization of the deepest Cambrian
158
layers (Ruiz et al., 2008) and of the basement, as shown by the K-Ar (Bonhomme and
159
Hassenforder, 1985) and zircon fission-track ages (Sebti et al., 2009) at about 330 Ma obtained
160
from various western Anti-Atlas basement samples. The Variscan collision developed a typical
161
thick-skinned tectonics in the Anti-Atlas (Burkhard et al., 2006). It was followed by a Late
162
Pennsylvanian-Permian period of erosion recorded by clastic deposits in the adjoining Tindouf
163
and Bechar basins (Conrad, 1972; Cavaroc et al., 1976) and responsible for most of the
164
exhumation of the Precambrian antiformal inliers.
165
The main Triassic-Liassic records in the Anti-Atlas domain consist of widespread dykes and
166
sills of the Central Atlantic Magmatic Province (CAMP; Hailwood and Mitchel, 1971; Hollard,
167
1973; Leblanc, 1973; Youbi et al., 2003) well dated in varied localities between 204±3 and 197±2
168
Ma (Ar-Ar plateau ages; Sebai et al., 1991). Triassic red beds are only preserved in the Siroua
169
Plateau next to the SAF (Chevallier et al., 2001; El Arabi et al., 2003). The Anti-Atlas was likely
170
uplifted and eroded during the Jurassic as a rift shoulder for both the Central Atlantic and the
171
Atlas Tethys Gulf. After a shallow burial during the Cretaceous-Eocene, recorded by scarce
172
Cretaceous outcrops onto the Siroua Plateau, the Anti-Atlas was definitely exhumed during the
173
Neogene, contemporaneously with the High Atlas.
174 175
3. Late Ediacaran extensional faulting 176
177
The timing and tectonic setting of the Ouarzazate Group have been, and still is the object
178
of active research. Several works have underlined the strong control of extensional faults on
179
clastic sedimentation and coeval volcanism in a continental rift environment (Choubert, 1963;
180
Azizi et al., 1990; Rjimati et al., 1992; Piqué et al., 1995; 1999; Thomas et al., 2002; Soulaimani
181
et al., 2003). Simultaneously, the magmatism shows an important chemical variation from bottom
182
to top (see Ezzouhairi et al., this volume). The calc-alkaline, arc-related bimodal magmatism that
183
dominates since the earliest post-orogenic periods (Saghro and Bou Salda Groups; see Liégeois et
184
al., this volume) evolves progressively during the late Ediacaran (Ouarzazate Group) to
185
widespread continental tholeiitic volcanism, then to alkaline magmatism in the Lower Cambrian
186
carbonates (Jbel Boho) at the very beginning of the Cambrian transgression. The late Ediacaran
187
period corresponds, according to most authors, to the collapse of the Pan-African chain and onset
188
of a transtensional-extensional regime subsequent to the Early Ediacaran transpressional regime
189
(Gasquet et al., 2008, and references therein). In the present section, we aim at featuring a
190
synthetic map of the main active faults during the late Ediacaran (Fig. 3), based on our field
191
works and the literature. We first document this map from west to east, with emphasis on the
192
basement faults that operated at the boundary of the future Anti-Atlas inliers, and then we
193
propose its broad interpretation.
194 195
3.1. Late Ediacaran faults in western Anti-Atlas
196
197
In western Anti-Atlas, the Ouarzazate Group clastic units show impressive and sudden
198
thickness changes (from 0 up to 800 m-thick) beneath the Lower Cambrian carbonates, suggesting
199
the role of synsedimentary normal faults. In several cases, the coarse, chaotic facies of the Ediacaran
200
conglomerates reveals the closeness of steep, fault related reliefs.
201 202
3.1.1. Paleofaults in the northern Kerdous
203
204
Outstanding examples of Late Ediacaran basement faults can be observed at the north-
205
eastern border of the Jbel Lkest (Fig. 3) (Soulaimani et al., 2004). Next to the Ida Ougnidif
206
locality (Fig. 4B), the Jbel Lkest Neoproterozoic metaquartzites are bounded by a set of NW-
207
trending faults running at the border of Upper Ediacaran clastic deposits topped by continental
208
tholeiites (Soulaimani et al., 2004; Soulaimani and Ouanaimi, 2011). Here, the uplifted quartzites
209
display various extensional structures such as tension gashes and mini-grabens whose orientation
210
is consistent with a normal throw along the main faults. Within the collapsed block to the NE, the
211
Ouarzazate Group conglomerates show poorly-sorted, angular quartzite clasts, implying proximal
212
deposition. Their thickness varies from a few tens of meters to several hundred meters beneath
213
the unconformable carbonates of the Adoudou Fm, which may results of late Ediacaran,
214
synsedimentary faulting, or post-sedimentary faulting prior to the Cambrian sedimentation, or
215
both. The occurrence of tens of meters of continental tholeiitic basalts on top of the
216
conglomerates at Ida Ougnidif, and that of similar tholeiites in several other places in the Anti-
217
Atlas (Youbi, 1998; Algouti et al., 2002; Soulaimani et al., 2004) supports the idea of an incipient
218
rifting event by the end of late Ediacaran.
219
The role of synsedimentary faulting is also made clear in the north of the Jbel Lkest massif,
220
where the destruction of the quartzites generated breccias and conglomerates collected in
221
topographic lows most likely associated with fault zones (O’Connor’s et al., 2010). Maximum
222
quartzite clasts size is commonly in excess of 1 m with frequently sub-angular shape. The occasional
223
presence of gabbro and granite clasts can be observed. The presence of interbedded tuffaceous
224
deposits in the conglomerates implies that explosive volcanism occurred during deposition. Indeed,
225
the J. Lkest example allows us confirming the importance of both synsedimentary and post-
226
sedimentary normal faulting along NW- to NNW-trending faults during the late Ediacaran, prior to
227
the Lower Cambrian transgression. It is worth noting that this fault corridor was the locus of
228
superimposed reactivations, i) during the Lower Cambrian, as shown by the west-ward directed
229
slumps and hydroplastic minor faults in the adjacent Tata Group limestones (Fig. 5F); and ii) during
230
the Variscan episode, with the development of N150E sub-vertical pressure solution cleavage in the
231
conglomerates along the Ida Ougnidif fault zone (Soulaimani and Ouanaimi, 2011). However, the
232
(moderate) reactivation of the paleofault does not hamper its identification as a late Ediacaran
233
normal fault.
234 235
3.1.2. Other basement faults in western Anti-Atlas
236
237
In the southern Kerdous massif, the Ouarzazate Group outcrops are relatively thin (Fig.
238
3). Clastic deposits and volcanic flows are controlled by NE-SW to E-striking faults, which
239
participate to the uplift of the basement (Chèvremont et al., 2005; Roger et al., 2005). The
240
ensuing horst and graben architecture has been invaded and sealed by marine limestones of the
241
Taroudant and Tata groups before being reactivated by the Variscan compression.
242
The Lakhssas synclinorium between the Kerdous and Ifni inliers offers another example of
243
Variscan inversion of upper Ediacaran horst and graben structures (Soulaimani and Bouabdelli,
244
2005). Gravimetric and magnetic data suggest the presence of an uplifted basement horst at depth
245
beneath the folded Cambrian limestones (Jbel Inter) in the axis of the synclinorium (Soulaimani,
246
1998) (Fig. 4C). An upper Ediacaran hemigraben can be restored south of the Tazeroualt
247
(southern Kerdous) horst (Fig. 3). The Variscan inversion mainly affected the western part of this
248
transect.
249
A similar configuration can be reported in the Tagragra of Akka inlier where Ediacaran
250
paleofaults are related to mylonitic zones in the Paleoproterozoic basement (Gasquet et al., 2001).
251
This is illustrated at the south-western tip of the inlier where a NW-dipping, N50E-striking
252
normal fault separates the Ouarzazate Group deposits in the hanging-wall from the
253
Paleoproterozoic basement in the footwall. This fault bounds a half-graben structure filled with
254
poorly-sorted conglomerates, which display chaotic facies with basement clasts up to 2 meters in
255
size close to the fault. The corresponding half-graben structure is sealed by the Adoudou Fm
256
carbonates.
257 258
3.2. Late Ediacaran structures in central Anti-Atlas
259
Extensional structures of late Ediacaran age can be recognized on both sides of the Anti-
260
Atlas Major Fault (AAMF), on the one hand in the Agadir Melloul and Iguerda inliers and on the
261
other hand in the Siroua and Bou Azzer inliers, south and north of the AAMF, respectively.
262 263
3.2.1. Agadir Melloul and Iguerda inliers
264
265
In the Agadir Melloul inlier, the Adrar Iguiguil hill consists of several hundred meters-
266
thick lower Neoproterozoic quartzites (Taghdout-Lkest Group), transgressive onto the
267
Paleoproterozoic basement and intruded by Neoproterozoic mafic dykes (Faure-Muret et al.,
268
1992; Soulaimani et al., 2013). The quartzite slab has been converted by gravity faults into a
269
system of roughly concentric tilted blocks. These blocks are covered by reddish, Ouarzazate
270
Group deposits including chaotic fault scarp breccias with quartzite clasts up to 2 m in size (Fig.
271
5B) and poorly-sorted conglomerates with dominant quartzite clasts and occasional
272
Paleoproterozoic clasts (Soulaimani et al., 2013).
273
In the eastern side of the Iguerda inlier (Fig. 3), the Aguinane Valley offers a perfect
274
illustration of upper Ediacaran extensional structures, which controls the Ouarzazate Group
275
deposits (Fig. 5C). On both sides of the valley, and particularly in its northern flank, a succession
276
of NE-striking listric faults and tilted basement blocks can be observed at the border of the inlier.
277
Volcaniclastic deposits of the Ouarzazate Group constitute the infilling of these half-graben
278
structures, which are topped and sealed by the Taroudant Group marine limestones (Hafid et al.,
279
2013). Therefore, the Iguerda basement, like many other Anti-Atlas inliers, was already a raised
280
block during the Ediacaran crustal extension.
281 282
3.2.2. Siroua and Bou Azzer inliers
283
284
The Ouarzazate Group formations outcrop widely north of the Anti-Atlas Major Fault
285
(AAMF) in the Siroua and Bou Azzer inliers. In the Siroua massif, the thick deposits are the
286
result of both explosive volcanic activity and rapid clastic sedimentation, controlled by normal
287
faults. Thomas et al. (2002) pointed out the importance of basement fractures in the onset and
288
evolution of the Ouarzazate Group basins there. Remarkably, the contents and thicknesses of
289
these basins are different north and south of the AAMF system. Whereas the group is dominated
290
by basement-derived conglomerates in the south, its succession is characterized in the north by a
291
major thickness of mainly acid volcanic/volcanoclastic rocks associated with a lower amount of
292
clastic sediments (Thomas et al. 2002). Therefore, the Ediacaran faulting clearly reactivated the
293
inherited AAMF in this area. It is not clear however whether the E-W fault-related grabens
294
developed by pure extension or by transtension.
295
In the Bou Azzer inlier, the Ouarzazate Group is only represented by its middle and upper
296
formations (Choubert, 1963; Boyer and Leblanc 1977). The volcaniclastic sequences overlie
297
unconformably the Pan-African ophiolites and oceanic arc units. NE-SW directed grabens and
298
half-grabens are observed beneath the unconformable carbonates of the Adoudou Fm (Fig. 5E).
299
According to Azizi-Samir et al. (1990), they correspond to en echelon faults related to the
300
sinistral reactivation of the WNW-striking major fault (AAMF) inherited from the Pan-African
301
orogeny. Coupled with the presence of the huge Ediacaran magmatism, an intense hydrothermal
302
activity developed during the extensional event, playing a major role in the formation of the well-
303
known ore deposits (Co-Ni-As-Ag-Au) of the area (Leblanc, 1975; Gasquet et al., 2005; Maacha
304
et al., 2011).
305 306
3.4. Late Ediacaran faults in eastern Anti-Atlas
307
Due to the gentle eastward plunge of the Anti-Atlas axis, the eastern Precambrian inliers
308
(Jbel Saghro and Ougnat; Fig. 3) display essentially upper Ediacaran volcanics and volcanoclastic
309
sequences of the Ouarzazate Group (Hindermeyer, 1954; Choubert, 1959; Choubert and Faure-
310
Muret, 1977; Paile, 1983; Ouguir et al., 1996; Abia et al., 2003; Gasquet et al., 2005; Raddi et al.,
311
2006a, b). These voluminous sequences unconformably overlie the lower Ediacaran folded series
312
(Saghro Group) that outcrop in small windows (Sidi Flah, Boumalne and Imiter in the Saghro
313
massif; Mellab and Ouin Oufrouh in the Ougnat massif). The Ouarzazate Group in both massifs is
314
also associated with important felsic to intermediate dyke swarms, gabbro stock (Raddi et al.,
315
2006a, b) and high-K pink granite intrusions (Fauvelet and Hindermeyer, 1951).
316 317
3.4.1. Saghro Massif
318
319
The Saghro massif is affected by several ENE- to NE-striking, tens of kilometres-long
320
faults (Fig. 3). These regional faults crosscut all the Precambrian rocks and locally the folded
321
Paleozoic series or even the Mesozoic-Cenozoic cover, which attests for repeated reactivation
322
events. This longitudinal fault system was active during the late Ediacaran as it controls the
323
thickness of the Ouarzazate Group deposits and the associated magmatic intrusions (Rjimati et al.,
324
1992; Walsh et al., 2012). As an example, the Tagmout graben, in the southern flank of the
325
Saghro massif, is an Ouarzazate Group-filled graben bounded by N60-striking faults associated
326
with the tholeiitic, 563 ±5 Ma-old Tagmout gabbro (Benziane et al., 2008). This fault system
327
shows frequent markers of sinistral strike-slip movements at the map scale. For instance, in the
328
Sidi Flah window, the N60-striking fault is accompanied by strike-slip duplexes and its left lateral
329
movement is accommodated by a horsetail at its eastern tip (Walsh et al., 2012). In addition, the
330
bulk left-lateral kinematics is consistent with the N10 trend of the oblique dikes, particularly of
331
those of the “Zone des dykes” locally dated at 564 ±7 (Bouskour mining district; Walsh et al.,
332
2008). In the north-eastern part of the Saghro massif, NE to ENE sinistral strike slip faults are
333
also reported as inherited from late Ediacaran paleofaults (El Boukhari et al., 2007, Malusà et al.,
334
2007; Massironi et al., 2007). These faults are likely contemporaneous to the E-trending fault
335
system described in the Imiter area (Ouguir et al., 1996; Cheilletz et al., 2002; Levresse et al.,
336
2004). The latter faults, inherited from old Pan-African dextral faults, were reactivated first as
337
normal faults controlling the Ouarzazate Group deposits, then as sinistral strike-slip fault (Ouguir
338
et al., 1994). These fractures served as hydrothermal drains leading to the first class Imiter Ag-Hg
339
deposits (Gaouzi et al., 2011).
340
NW-striking faults were also active during the late Ediacaran in the Saghro massif. This is
341
documented in the Boumalne area where N120E, kilometre-long normal faults control the Jbel
342
Habab graben filled up with Ouarzazate Group formations (El Boukhari et al., 2007). The bordering
343
faults of this graben are sealed in the northwest by Cambrian beds, thus excluding significant post-
344
Precambrian reactivations.
345
3.4.2. Ougnat Massif
346
347
In the easternmost Anti-Atlas, i.e. the Jbel Ougnat massif, the Ouarzazate Group is
348
dominated by felsic volcanics, mainly ignimbritic sheets, associated with various magmatic
349
intrusions, either gabbroic or granitic (Paile 1983; Abia et al., 2003). The scarcity of clastic
350
deposits belonging to the group hampers defining synsedimentary grabens or half-grabens. In
351
contrast, the widespread dolerite and granite dykes can be used as indicators of extension coeval
352
with magmatism. In most areas, the mafic dykes strike dominantly NE-SW (Fig. 3; Raddi et al.
353
2006a, b), crosscutting the Saghro Group basement as well as the Ouarzazate Group volcanics
354
and plutons. A great number of N-S and NW-SE mafic or felsic dykes also occur in the western
355
and central areas. Around the Bou Madine mine, the NW-trending felsic dykes are crosscut by the
356
N-striking mafic ones. In this area, major sinistral N30-striking faults are associated with N160-
357
striking tension joints hosting the epithermal polymetallic ore deposit (Abia et al., 2003). As a
358
whole, the dyke orientation suggests a multidirectional extension during the upper Ediacaran
359
magmatic evolution, with a dominant NW-SE direction of extension. Contrary to the Jbel Saghro,
360
it is not clear here whether the Ouarzazate Group was controlled by a global left-lateral
361
transtensional context along N70 faults along the north and south borders of the massif.
362 363
4. Paleozoic evolution and inherited basement faults
364
365
4.1. Early Cambrian rifting 366
367
In many places, as reported above, the base of the Adoudounian carbonates conformably
368
overlie the Ouarzazate Group without apparent stratigraphical gap (e. g. eastern flank of the Kerdous
369
inlier (Fig. 5A). This is verified mainly in the western and southern part of the Anti-Atlas. However,
370
in other places, this transition is marked by an angular unconformity on top of the tilted formations
371
of the Ouarzazate Group (e. g. NE of Taliwine, Fig. 5D). Everywhere, extensional tectonics went on
372
during the Early Cambrian, as illustrated hereafter.
373
The Cambrian marine transgression was progressive from the western “Gulf of Souss” (Choubert
374
and Marçais, 1952) to the central and eastern Anti-Atlas (Destombes et al., 1985). The Adoudou
375
Fm-Lower Cambrian thickness, up to 3000 m in the western Anti-Atlas, decreases gradually
376
eastward while continental influences increase. The role of extensional tectonics during the
377
evolution of the Lower Cambrian basin has been emphasized either by structural (Soulaimani et al.,
378
2003, Soulaimani and Piqué, 2004, Gasquet et al., 2005) or sedimentological studies (Benssaou and
379
Hamoumi, 2003; Chbani et al., 1999; Algouti et al., 2002). Thickness and facies changes have been
380
used to map a NE-striking Lower Cambrian graben (Benssaou and Hamoumi, 2003). The Adoudou
381
limestones often show disruptions by synsedimentary extensional faults and slump structures
382
(Soulaimani et al., 2003) (Fig. 5G, H). Large scale synsedimentary normal faults, detachments and
383
slump folds can be observed south of Tiouine (westernmost Saghro massif; Fig. 5E) in the lower
384
Adoudou clastics (“Série de base”) and limestones (“Calcaires inférieurs”) and the Taliwine Fm
385
(“Série lie-de-vin”), suggesting the occurrence of tilted basement block underneath (Saddiqi et al.,
386
2011).
387
Lower Cambrian faults are also exposed in the Issafen syncline (Fig. 3) where many NNE-
388
striking, kilometre scale faults cut preferential levels of the Lower Cambrian units suggesting a
389
synsedimentary deformation. Likewise, alkaline volcanism went on during the Early Cambrian as
390
a continuation of the tholeiitic flows at the top of the Ouarzazate Group (sect. 3.1.1). The most
391
important record corresponds to the J. Boho volcano whose flows are interleaved in the
392
uppermost Adoudou beds at the southern side of the Bou Azzer inlier (Ducrot and Lancelot,
393
1977; Leblanc and Lancelot, 1980; Alvaro et al., 2006). This stratigraphic position is consistent
394
with the 529 ± 3 Ma U-Pb age yielded by the J. Boho syenite (Gasquet et al., 2005).
395
In the Anti-Atlas, Early Cambrian rifting aborted before the widespread sandy sedimentation of
396
the Asrir Fm (former “Grès terminaux”) now attributed to the base of Middle Cambrian (Geyer and
397
Landing, 1995, 2004). The thickness of the Lower Cambrian and Asrir formations decreases
398
eastward (Destombes et al., 1985; Buggisch and Siegert 1988) and they vanish finally at the
399
northern border of the eastern Saghro and Ougnat inliers (Fig. 6A). Abrupt thickness changes clearly
400
suggest the activity of (inherited) basement faults with N70 and N110 dominant directions. Minor
401
faults with NW-directed normal throw are illustrated at the south border of the Precambrian inlier
402
(Fig. 6B).
403
During the Middle Cambrian transgression, sedimentation is dominated by fairly continuous,
404
mainly south-derived detrital input with minor disconformities. Extensional tectonics is evidenced
405
by mega-slumps and seismites in eastern Anti-Atlas (Raddi et al., 2007), associated with a
406
voluminous mafic volcanism (Destombes, 2006c, d; Raddi, 2014). This volcanism is dominated
407
by trachy-basalt flows and microdolerite stocks and dykes with alkaline affinity. Distribution of
408
the paleovolcanoes seems mostly controlled by N70- and N120-striking faults.
409 410
The Upper Cambrian (Furongian) corresponds to the most important interruption in the
411
Palaeozoic sedimentation all over Morocco except in restricted areas of western Anti-Atlas
412
(Destombes and Feist, 1987), western High Atlas (Cornée et al., 1987) and western Meseta
413
(André et al., 1987; El Attari et al., 1997; Mergl et al., 1998). The scarcity of stratigraphic records
414
has not been explained yet; it could result from an uplift of central and eastern Morocco as a
415
tectonic shoulder of the spreading Iapetus.
416 417
4.2. Ordovician-Lower Devonian, the subsiding platform
418
419
The Lower and Middle Ordovician deposits consist of several hundred meters-thick
420
pelites and sandstones with frequent hiatus, disconformities and shallow water ferruginous
421
oolithes (Destombes et al., 1985; Destombes, 2006a-d; Marante, 2008). The maximum of
422
subsidence of the Ordovician Saharan platform is observed in the western and central Anti-Atlas.
423
Contrastingly, the eastern Anti-Atlas was poorly subsident, with a complete hiatus of the
424
Tremadoc deposits around the Ougnat massif. Isopachs are dominantly E-W (i.e. parallel to the
425
former coast line concealed beneath the Tindouf Basin) in most of the belt up to the Upper
426
Ordovician (Caradoc), when they turn to NW-SE in eastern Anti-Atlas (Destombes et al., 1985).
427
This suggests the first activation of basement faults along the Ougarta Belt trend and its
428
northward continuation, namely the Ougnat-Ouzina axis. At the scale of the entire Anti-Atlas,
429
Marante (2008) shows that the Pan-African suture zone was reactivated into a crustal flexure zone
430
during the Middle Ordovician.
431
The NW-SE direction is also that of the Hirnantian tilloids and Rhuddanian black shales along
432
the paleofjords of eastern Anti-Atlas (Destombes, 2006d; Le Heron, 2007). However, after this
433
period of fault activity, the Saharan platform entered again a period of quiescent, although uneven
434
subsidence that lagged during most of the Silurian and Lower Devonian.
435 436
4.3 Middle-Late Devonian to Early Carboniferous: marginal platform dislocation
437
438
During the Middle-Late Devonian the Saharan platform was converted into a complex of uplifted
439
blocks (minor platforms) and downthrown basins (Fig. 2). This “disintegration” event (Wendt,
440
1985) is well documented in the eastern Anti-Atlas (Wendt and Belka, 1991; Baidder, 2007;
441
Baidder et al., 2008). There, the Devonian normal fault pattern indicates a multi-directional
442
extension with a dominant northward direction (Fig. 7). The most important faults are inherited
443
from the NNW- and ENE-trending basement faults active during the Precambrian. This is clearly
444
the case with the Oumjerane-Taouz Fault, which is the eastern continuation of the Pan-African
445
main fault zone (AAMF). In the south of the Ougnat massif, the N-Mecissi fault is another
446
example of ENE-striking, N-dipping Late Devonian faults (Raddi et al., 2007). The extensional
447
reactivation of the NNW-striking faults on both sides of the Ougnat-Ouzina axis determines the
448
differentiation of two subsiding basins, namely the Maider and South Tafilalt basins, bounded by
449
shallow pelagic platforms (Hollard 1974, 1981; Wendt 1985,1988; Baidder et al., 2008).
450
In the Western Anti-Atlas, faulting and paleogeographic differentiation occurred earlier, as
451
shown by the important sequence variations in the Lower to Middle Devonian Rich Group
452
(Ouanaimi and Lazreq, 2008), contrasting with the monotonous basinal facies of the Upper
453
Devonian successions.
454
Finally, by the end of the pre-orogenic evolution, the Lower Carboniferous sedimentation was
455
controlled by the same basement fault pattern as during the Upper Devonian. The activity of the
456
Oumjerane-Taouz fault zone is documented by coarse conglomerates, slumpings and
457
olistostromes at the south border of the Tafilalt Basin, next to the shallow platform further in the
458
south (Jebel Begaa carbonates; Hollard, 1970). Likewise, turbidites and olistostromes
459
accumulated along the northern boundary of the eastern Anti-Atlas (Tineghir area north of Saghro
460
and Ougnat inliers) and adjacent Bechar Basin (Ben Zireg), in relation with ENE-striking normal
461
fault activation (Michard et al., 1982; Soualhine et al., 2003; Cerrina Feroni et al., 2010).
462
The geodynamic framework of the Middle-Late Devonian-Early Carboniferous sedimentation
463
has been discussed by Frizon de Lamotte et al. (2013) from North Africa to Arabia. They point to
464
a major thermo-mechanical event at the scale of northern Gondwana resulting in diffuse
465
extensional deformation (rifting) with contrasting uplifted archs and deep basins. This can be
466
regarded as the consequence of the Laurussia plate incipient subduction beneath Gondwana along
467
the nascent Variscan Belt.
468 469
5. Variscan inversion of the inherited structures
470
471
5.1. Faulted inliers of the foreland fold belt: General
472
473
During the Variscan Laurussia-Gondwana collision, the Paleozoic series of the Anti-Atlas
474
basin has been folded whereas large blocks of its Proterozoic basement were uplifted as
475
antiformal inliers or “boutonnières” (Fig. 1 and 7). Except the westernmost part of the belt where
476
narrow NNE- trending Cambrian ridges along the Atlantic coast are affected by a west-verging
477
(craton-ward) thin-skinned thrust system (Soulaimani, 1998; Belfoul et al., 2001), most of the
478
Anti-Atlas shows an obvious implication of the Precambrian basement defining a thick-skinned
479
tectonic style (Hassenforder, 1987; Soulaimani et al., 1997; Belfoul et al. 2001; Caritg et al.,
480
2004; Helg et al. 2004; Burkhard et al., 2006; Baidder et al., 2007; Raddi et al., 2007; Soulaimani
481
and Burkhard, 2008; Michard et al., 2008, 2010).
482
In the basement samples collected throughout western and central Anti-Atlas, zircon fission-
483
track (ZFT) ages cluster between 340 ± 20 and 306 ± 20 Ma (average age 321 ± 21 Ma; Sebti et
484
al., 2009; Oukassou et al., 2013). These ZFT ages are consistent with the K/Ar and 40Ar-39Ar
485
results from the Kerdous area (Bonhomme and Hassenforder, 1985; Soulaimani and Piqué, 2004).
486
Therefore a single thermal event has been responsible for resetting of the various dating systems
487
about 310-330 My ago (late Visean-Bashkirian), being followed by rapid cooling below 240 ±
488
20°C (closure temperature for zircon fission-track dating). Peak temperature in the upper part of
489
the basement hardly exceeded T=300±20°C according to the epizonal recrystallization of the
490
overlying Cambrian rocks of the area (Ruiz et al., 2008). As a consequence of the low
491
temperature of the basement rocks during collision, they were deformed in brittle conditions, so
492
as the Precambrian inliers correspond to strongly faulted basement-cored uplifts of various size
493
and orientation (Fig. 1). In many cases, they are bounded at least on one side by steeply-dipping
494
faults either exposed or concealed under the Adoudounian-Cambrian beds (Fig. 7). Along other
495
sides, they are overlain by the folded Paleozoic series detached on numerous décollements levels
496
linked to the main incompetent formations. In the western Anti-Atlas, the deepest décollement
497
horizon is located in the Lower Cambrian Taliwine Fm (“Lie-de-Vin” pelites), whereas in the
498
eastern Anti-Atlas it is located in the Middle Cambrian Internal Feijas Gp (“Schistes à
499
Paradoxides” pelites). Evidence for the ancestry of most bounding-faults of the Anti-Atlas inliers
500
and their Variscan reactivation has been reported in many areas of the western Anti-Atlas
501
(Hassenforder, 1987; Piqué et al. 1987; Soulaimani 1998). In the following, we investigate key
502
examples from the whole belt attesting for the Variscan reactivation of late Ediacaran structures
503
with various rates and motions.
504 505
5.2. Totally inverted listric paleofaults
506
507
Total inversion of former listric faults can be only observed in the westernmost part of the
508
belt, close to the front of the Mauritanides thrusts (Fig. 1). The Bas Draa inlier (Bourcart, 1937;
509
Choubert and Faure-Muret, 1969), here discussed is bounded along its south-eastern side by
510
steeply-dipping reverse faults (Fig. 8A, B; Soulaimani et al., 1997). During the Variscan collision,
511
the basement block was simultaneously uplifted and thrust southeast-ward inducing folding and
512
axial-plane cleavage development in the sedimentary cover at its front. The south-eastern border
513
of the inlier corresponds to one of the Ediacaran-Cambrian faults that bounded the western Anti-
514
Atlas Cambrian rift. Thus, in the Bas Draa case study, the Variscan contraction brought the
515
hanging-wall of the Cambrian paleofault higher than the footwall removing entirely the
516
extensional geometry (Fig. 8C).
517 518
5.3. Incompletely inverted paleofaults
519
520
Contrary to the Bas Draa example, many of the basement faults bounding the Anti-Atlas
521
inliers still preserve normal throw. The Assaragh fault is an outstanding example of such
522
incompletely inverted paleofault. This fault constitutes the eastern branch of NNE-SSW to N-S
523
fault system that affects the Agadir Melloul-Assaragh basement (Fig. 9A). It separates the Iguerda
524
Paleoproterozoic inlier from the Aguinane inlier. The latter is characterized by the double
525
unconformity of the Adoudou Fm limestones onto the Ouarzazate Group volcaniclastics and the
526
Paleoproterozoic schists. The present throw of the Adoudou unconformity between both inliers
527
looks like a normal fault throw, but mapping of the fault zone shows that both the Ediacaran lavas
528
and Adoudou limestones are folded with geometry only compatible with a reverse basement fault
529
underneath (Fig. 9B-C). The Assaragh fault then appears as a polyphase fault that acted as a
530
normal fault during and/or after the late Ediacaran (rifting phases, see sect. 3 and 4.1), and as a
531
reverse fault during the Variscan compression, but with a lesser throw than the previous normal
532
one. It is worth noting that the Assaragh fault parallels the nearby Paleoproterozoic Lamdint shear
533
zones (Faure-Muret et al., 1992, Hafid et al., 2013), suggesting that the paleofault itself is
534
inherited from a much older weakness zone.
535 536
5.4. Blind inverted paleofaults
537
538
Outside the Anti-Atlas axis where inverted basement faults are exposed, many other
539
basement faults are buried beneath the Paleozoic cover. In the absence of convenient seismic
540
data, these hidden basement faults are revealed by the propagation of secondary faults cutting
541
through the entire Palaeozoic cover and demonstrating a normal, then reverse or reverse strike-
542
slip activity.
543
In the western Anti-Atlas, these faults are organized in two main directions, broadly E-W and
544
N-S. The most important E-W lineament is the Anti-Atlas Major Fault (AAMF; sect. 3.2) that
545
crosscuts the Jebel Bani and extends eastward in the Zagoura and Oumjerane-Taouz faults (Fig.
546
7). This is a Pan-African compressional structure (sect. 2) reactivated during the Ediacaran and
547
the Paleozoic as an extensional fault, and again as a reverse-transcurrent faults during the
548
Variscan compression.
549
The E-W Tata fault (Fig. 7) parallels the AAMF in the south and has a strong control on the
550
Variscan structures (Hassenforder, 1987; Caritg et al., 2004). Other remobilized basement faults
551
do not broke the surface and are only expressed by fault-propagation folds. South of the Tata
552
Fault, the Adrar Zouggar-Addana Ordovician anticlinorium overlies a hidden Precambrian high
553
(Michard, 1976; Burkhard et al., 2006). Still more in the southwest, the N70°E-striking shear
554
zone between the Anti-Atlas and the cratonic Tindouf Basin (Fig. 7) is another conspicuous
555
example of basement paleofault whose inversion induced en echelon folds within the Jebel Rich
556
Devonian sequences in response to the dominating SE-directed compression (Michard 1976;
557
Jeannette and Piqué 1981; Soulaimani et al., 1997; Michard et al., 2010).
558
Between the AAMF and Tata E-W basement faults, an N-S fault system is represented by the
559
Lakhssas Plateau shear zone and the Agadir Melloul lineament as principal examples. The
560
Lakhssas Plateau shear zone (Fig. 4C) is the expression in the Lower Cambrian of the inversion
561
of a listric basement fault carrying the Ifni block eastward against the Kerdous (Soulaimani and
562
Bouabdelli, 2005; Michard et al., 2010). The Agadir Melloul lineament is N-S polyphase fault
563
system that exhumed the granitic basement against the Ouarzazate Group before the Adoudou Fm
564
deposition. The Assaragh branch of the lineaments still a normal fault, as described above (§ 5.3,
565
Fig. 9),but laterally to the south, the fault system is expressed in the Adoudounian-Lower
566
Cambrian cover as an impressive fold system (Soulaimani et al., 2013).
567 568
In the eastern Anti-Atlas, both NW-SE and E-W basement paleofaults reactivated as reverse
569
strike-slip faults can be observed. In particular, two major NW-striking faults occur on both sides
570
of the Ougnat-Ouzina Cambrian-Ordovician axis. They operated as the boundaries of a mega-
571
dextral shear zone in response to the N-S compression that first prevailed there (Baidder et al.,
572
2008).
573 574
575
6. Reactivation of basement faults during the post-Variscan Anti-Atlas evolution
576
577
The Mesozoic-Cenozoic tectonics of the Anti-Atlas has long been misunderstood or at
578
least underestimated, probably due to the lack of related deposits upon most of the chain (sect. 2).
579
The Anti-Atlas was erroneously considered as totally outside the Alpine Atlas system (Choubert
580
and Marçais, 1952; Michard, 1976). Contrastingly and mostly based on apatite fission-track
581
(AFT) data, it is now accepted that the Anti-Atlas underwent superimposed phases of post-
582
Variscan burial and exhumation (Missenard et al., 2006; Malusà et al., 2007; Ruiz et al., 2010;
583
Oukassou et al., 2013). In the following, we examine the role of the inherited basement faults
584
during three significant periods, i.e. the Triassic-Jurassic, the Cretaceous-Eocene and the Neogene
585
to Present.
586 587
6.1. Triassic rifting and Jurassic emersion
588
589
The Triassic-Early Jurassic rifting of Pangea culminated with the CAMP magmatic event at
590
ca. 200 Ma (Sebai et al., 1991; Verati et al., 2007). At that time, the Anti-Atlas area basically
591
belongs to the uplifted shoulder of the rift zone, which extended in the future Atlantic margin and
592
High Atlas basin. The shoulder erosion sourced the redbeds of the Kenadza-Bechar Basin in the east
593
(Fabre, 2005), those of the High Atlas in the north (Beauchamp et al., 1996, 1999; El Arabi et al.,
594
2006; Frizon de Lamotte et al., 2009) and those of the Atlantic margin onshore and offshore
595
(Mustaphi et al., 1997; Hafid et al., 2006).
596
However, the stability of the Anti-Atlas during the latest Triassic times is a matter of debate. Late
597
Triassic extension certainly affected the Variscan Anti-Atlas belt, being recorded by the intrusion of
598
several NE-trending dykes and associated sills of gabbro and dolerite (Hailwood and Mitchell, 1971;
599
Hollard, 1973; Sebai et al., 1991) (Fig. 10). Such large intrusions strongly suggest that basaltic
600
trapps covered the whole Anti-Atlas 200 Ma ago, associated with some continental red beds as those
601
observed on the northern edge of the Siroua Plateau (Chevallier et al., 2001). Besides of these dykes,
602
normal faults (probably inherited from Variscan ones) contributed to the lowering of the post-
603
Variscan peneplain at the northern border of the eroded belt (Robert-Charrue and Burkhard, 2008).
604
The faulted Anti-Atlas remained or became again a subaerial, eroded domain during the Jurassic,
605
as shown by the elimination of any Triassic basalts or sediments in most of the domain beneath the
606
unconformable Early Cretaceous continental deposits of the hamadas (Zouhri et al., 2008). Apatite
607
fission-track (AFT) studies on samples from the Precambrian inliers yielded varied, but mostly
608
Jurassic-Cretaceous apparent ages (Fig. 10). They are presented at the end of the section as they
609
heavily depend on the Cretaceous-Eocene evolution.
610 611
6.2. Cretaceous-Eocene burial
612
613
Continental red beds began to accumulate onto the eroded Anti-Atlas fold belt possibly as
614
early as the Barremian (like in the High Atlas and Meseta domains) or at least during the lower
615
Cenomanian (fossiliferous beds of the “Continental intercalaire” at the bottom of the Kem Kem and
616
hamada plateaus; Zouhri et al., 2008). Sedimentation went on in shallow marine conditions during
617
the Late Cretaceous (Cenomanian-Turonian carbonates, “Senonian” = Coniacian-Maastrichtian
618
gypsiferous marls) and the Lower-Middle Eocene (Swezey, 2009). Reactivation of paleofaults is
619
likely during the Early and “Middle” Cretaceous, which corresponds to the breakdown of Gondwana
620
and early drifting of the South Atlantic (Guiraud et al., 2005; Geraldes et al., 2013).
621 622
6.3. Oligocene to Present
623
624
The Cretaceous-Eocene deposits of the Anti-Atlas and Saharan hamadas are overlain by
625
unconformable Oligocene (?)-Neogene continental deposits that overlie directly the Anti-Atlas
626
axis in some areas (e.g. west of Ouarzazate or south-east of Erfoud). This echoes the stratigraphic
627
evolution of the Atlas System immediately in the north (Frizon de Lamotte et al., 2009).
628
The Cretaceous tabular plateaus all around the belt yield evidence of significant “Atlasic”
629
deformations, mainly faults, but also very open folds. A number of ENE-trending faults are
630
observed along the southern Sub-Atlas Zone (Souss and Ouarzazate Basins; Fig. 10). They
631
broadly parallel the SAF and correspond to inherited basement faults as the SAF itself. To the
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south of the belt, the Tata Fault is reported as a recently active fault with hundred meters of
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normal throw (Choubert and Marçais, 1952).
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In the western Anti-Atlas, N-striking faults control the present day topography of the area.
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Isolated Cretaceous outcrops preserved south and north of Tiznit (Waters et al., 2001) unravel the
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large throw (>400 m) of these normal faults, which likely correspond to the negative inversion of
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the Variscan faults beneath the Lakhssas Plateau. Likewise, the northern part of the Siroua
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Plateau is affected by a set of N-striking faults that bound an equal number of Cretaceous mini-
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grabens. This setting is consistent with the N-S directed push of the Marrakech High Atlas against
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the Siroua Plateau.
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In the eastern Anti-Atlas, post-Eocene, ENE-striking faults have been mapped along the
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northern Saghro and Ougnat inliers. Likewise, north of Erfoud, very open E-trending folds affect
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the Cenomanian-Turonian slab of the Meski Hamada that detached on the reddish Cenomanian
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clays (Saddiqi et al., 2011). Further in the south-east, the J. Zorg fault close to Taouz is oriented
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N140E, parallel to the Ougarta Belt (Fig. 10); its normal throw exceed the thickness of the
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neighbouring Cretaceous plateau (ca. 200 m-thick). The latter is crosscut by several ENE-striking
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faults, parallel to the Oumjerane-Taouz basement fault.
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Globally, the Cretaceous-Cenozoic cover of the Anti-Atlas dips north and south along the
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northern and southern borders of the mountain belt, respectively. Thus, the belt looks like a
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lithosphere-scale fold (“pli de fond”) coeval with the neighbouring High Atlas and related to the
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Alpine collision (Frizon de Lamotte et al., 2000, 2009). However, the relatively high relief of
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both the High Atlas and Anti-Atlas also relies on a hot mantle anomaly extending obliquely from
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the Ifni-Siroua area to the Middle Atlas and eastern Rif region (Missenard et al., 2006; Fullea et
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al., 2010 and references therein).
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Numerous neotectonic records such as faulted Holocene spring tuffs (Boudad et al., 2003;
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Weisrock et al., 2008) occur in the Anti-Atlas. Finally, seismicity is relatively important and
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reveals present-day fault reactivation events. The Rissani earthquake (1992) occurred in the core
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of the eastern Anti-Atlas (Hahou et al, 2003; Bensaïd et al., 2009). Its focal mechanism have been
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ascribed to dextral displacements along E-W directed strike-slip faults, parallel to the many
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Paleozoic faults of the area (e.g. Erfoud fault).
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6.4. Low-temperature geochronology
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Numerous apatite fission-track (AFT) and zircon U-Th/He (ZHe) studies helped
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reconstructing the vertical movements of the Anti-Atlas belt during the post-Variscan times. The
